Evolution of the Atmosphere

10/04/06

The evolution of the atmosphere
can be divided into four separate stages:

Origin
Chemical/pre-biological
Microbial era
Biological era

The composition
of the present atmosphere however required the formation of oxygen
to sufficient levels to sustain life, and required life to create
the sufficient levels of oxygen. This era of evolution of the atmosphere
is called the "Biological Era."

The Biological Era - The Formation of Atmospheric
Oxygen

The biological era was marked by the simultaneous decrease in atmospheric
carbon dioxide (CO2) and the increase in oxygen (O2) due to life
processes. We
need to understand how photosynthesis could have led to maintenance
of the ~20% present-day level of O2.

The build up of oxygen had three major consequences.

Firstly, Eukaryotic metabolism could only have begun once the level
of oxygen had built up to about 0.2%, or ~1% of its present abundance.
This must have occurred by ~2 billion years ago, according to the
fossil record. Thus, the eukaryotes came about as a consequence
of the long, steady, but less efficient earlier photosynthesis carried
out by
Prokaryotes.

Figure 1. Photolysis of water vapor and carbon
dioxide produce hydroxyl and atomic oxygen, respectively, that, in
turn, produce oxygen in small concentrations. This process produced
oxygen for the early atmosphere before photosynthesis became dominant.

Oxygen increased in stages, first through photolysis
(Figure 1) of water vapor and carbon dioxide by ultraviolet energy
and, possibly, lightning:

H2O -> H + OH

produces a hydroxyl radiacal (OH) and

CO2 -> CO+ O

produces an atomic oxygen (O). The OH is very reactive
and combines with the O

O + OH -> O2 + H

The hydrogen atoms formed in these reactions are light
and some small fraction excape to space allowing the O2 to build
to a very low concentration, probably yielded only about 1% of the
oxygen available today.

Secondly, once sufficient oxygen had accumulated in the stratosphere,
it was acted on by sunlight to form ozone, which allowed colonization
of the land. The first evidence for vascular plant colonization
of the land dates back to ~400 million years ago.

Thirdly, the availability of oxygen enabled a diversification of
metabolic pathways, leading to a great increase in efficiency. The
bulk of the oxygen formed once life began on the planet, principally
through the process of photosynthesis:

6CO2 + 6H2O <--> C6H12O6
+ 6O2

where carbon dioxide and water vapor, in the presence of light,
produce organics and oxygen. The reaction can go either way as in
the case of respiration or decay the organic matter takes up oxygen
to form carbon dioxide and water vapor.

Life started to have a major impact on the environment once photosynthetic
organisms evolved. These organisms fed off atmospheric carbon dioxide
and converted much of it into marine sediments consisting of the innumerable
shells and decomposed remnants of sea creatures.

Cumulative history of O2 by photosynthesis through
geologic time.

While photosynthetic life reduced the carbon dioxide content of the
atmosphere, it also started to produce oxygen. The oxygen did not
build up in the atmosphere for a long time, since it was absorbed
by rocks that could be easily oxidized (rusted). To this day, most
of the oxygen produced over time is locked up in the ancient "banded
rock" and "red bed" rock formations found in ancient
sedimentary rock. It was not until ~1 billion years ago that the reservoirs
of oxidizable rock became saturated and the free oxygen stayed in
the air. The figure illustrates a possible scenario.

We have briefly mentioned the difference between reducing (electron-rich)
and oxidizing (electron hungry) substances. Oxygen is the most important
example of the latter type of substance that led to the term oxidation
for the process of transferring electrons from reducing to oxidizing
materials. This consideration is important for our discussion of atmospheric
evolution, since the oxygen produced by early photosynthesis must
have readily combined with any available reducing substance. It did
not have far to look!

We have been able to outline the steps in the long drawn out process
of producing present-day levels of oxygen in the atmosphere. We refer
here to the geological evidence.

When the oceans first formed, the waters must have dissolved enormous
quantities of reducing iron ions, such as Fe2+. These ferrous
ions were the consequences of millions of years of rock weathering
in an anaerobic (oxygen-free) environment. The first oxygen produced
in the oceans by the early prokaryotic cells would have quickly been
taken up in oxidizing reactions with dissolved iron. This oceanic
oxidization reaction produces Ferric oxide Fe2O3
that would have deposited in ocean floor sediments. The earliest evidence
of this process dates back to the Banded Iron Formations, which reach
a peak occurrence in metamorphosed sedimentary rock at least 3.5 billion
years old. Most of the major economic deposits of iron ore are from
Banded Iron formations. These formations, were created as sediments
in ancient oceans and are found in rocks in the range 2 - 3.5 billion
years old. Very few banded iron formations have been found with more
recent dates, suggesting that the continued production of oxygen had
finally exhausted the capability of the dissolved iron ions reservoir.
At this point another process started to take up the available oxygen.

Red Beds

Once the ocean reservoir had been exhausted, the newly created oxygen
found another large reservoir - reduced minerals available on the barren
land. Oxidization of reduced minerals, such as pyrite FeS2
, exposed on land would transfer oxidized substances to rivers and out
to the oceans via river flow. Deposits of Fe2O3
that are found in alternating layers with other sediments of land origin
are known as Red Beds, and are found to date from 2.0 billion years
ago. The earliest occurrence of red beds is roughly simultaneous with
the disappearance of the banded iron formation, further evidence that
the oceans were cleared of reduced metals before O2 began
to diffuse into the atmosphere.

When the red bed reservoir became exhausted too
(although it is continually being regenerated through weathering)
and oxygen finally started to accumulate in the atmosphere itself.
This signal event initiated eukaryotic cell development, land
colonization, and species diversification. The oxygen built up to today's value only after the colonization
of land by green plants, leading to efficient and ubiquitous photosynthesis.
The current level of 20% seems stable.

The Oxygen Concentration Problem.

Why does present-day oxygen sit at 20%? This is not a trivial question
since significantly lower or higher levels would be damaging to life.
If we had < 15% oxygen, fires would not burn, yet at > 25% oxygen,
even wet organic matter would burn freely.

The Early Ultraviolet Problem

The genetic materials of cells (DNA) is highly susceptible to damage
by ultraviolet light at wavelengths near 0.25 µm. It is estimated
that typical contemporary microorganisms would be killed in a matter
of seconds if exposed to the full intensity of solar radiation at
these wavelength. Today, of course, such organisms are protected by
the atmospheric ozone layer that effectively absorbs light at these
short wavelengths, but what happened in the early Earth prior to the
significant production of atmospheric oxygen? There is no problem
for the original non-photosynthetic microorganisms that could quite
happily have lived in the deep ocean and in muds, well hidden from
sunlight. But for the early photosynthetic prokaryotes, it must have
been a matter of life and death.

It is a classical "chicken and egg" problem. In order to
become photosynthetic, early microorganisms must have had access to
sunlight, yet they must have also had protection against the UV radiation.
The oceans only provide limited protection. Since water does not absorb
very strongly in the ultraviolet a depth of several tens of meters
is needed for full UV protection. Perhaps the organisms used a protective
layer of the dead bodies of their brethren. Perhaps this is the origin
of the stromatolites - algal mats that would have provided adequate
protection for those organisms buried a few millimeters in. Perhaps
the early organisms had a protective UV-absorbing case made up of
disposable DNA - there is some intriguing evidence of unused modern
elaborate repair mechanisms that allow certain cells to repair moderate
UV damage to their DNA. However it was accomplished, we know that
natural selection worked in favor of the photosynthetic microorganisms,
leading to further diversification.

Fluctuations in Oxygen

The history of macroscopic life on Earth is divided into three great
eras: the Paleozoic, Mesozoic and Cenozoic. Each era is then divided
into periods. The latter half of the Paleozoic era, includes the Devonian
period, which ended about 360 million years ago, the Carboniferous
period, which ended about 280 million years ago, and the Permian period,
which ended about 250 million years ago.

According to recently developed geochemical models, oxygen levels
are believed to have climbed to a maximum of 35 percent and then dropped
to a low of 15 percent during a 120-million-year period that ended
in a mass extinction at the end of the Permian. Such a jump in oxygen
would have had dramatic biological consequences by enhancing diffusion-dependent
processes such as respiration, allowing insects such as dragonflies,
centipedes, scorpions and spiders to grow to very large sizes. Fossil
records indicate, for example, that one species of dragonfly had a
wing span of 2 1/2 feet.

Geochemical models indicate that near the close of the Paleozoic
era, during the Permian period, global atmospheric oxygen levels dropped
to about 15 percent, lower that the current atmospheric level of 21
percent. The Permian period is marked by one of the greatest extinctions
of both land and aquatic animals, including the giant dragonflies.
But it is not believed that the drop in oxygen played a significant
role in causing the extinction. Some creatures that became specially
adapted to living in an oxygen-rich environment, such as the large
flying insects and other giant arthropods, however, may have been
unable to survive when the oxygen atmosphere underwent dramatic change.

Structure of the Atmosphere

In large measure, the atmosphere has evolved in response to and controlled
by life processes. It continues to change as a consequence of human
activities, but at a rate that is far in excess of the rate of previous
evolutionary change. The atmosphere controls the climate and ultimately
determines the quality of life on Earth. We will begin our discussion
with a brief review of the composition and structure of the present-day
atmosphere. Then we will discuss the major events in the evolution
of the atmosphere that led to its current state. We will discuss some
important tools along the way that will prove useful in many settings.

The word "structure" is used in atmospheric physics
to mean the vertical profile of particular variables of interest
(such as temperature, density, pressure, etc.)

Atmospheric structure is subdivided into four thermal layers or
"-spheres" that are divided by transition regions or "-pauses".
The nomenclature used dates back to the 1950's and is based on the
measured temperature profile of the atmosphere. Figure 1 illustrates
the temperature profile and the names used for the different regions.

As discussed earlier, the ground heats up due to the absorption of
visible light from the Sun. The warm ground, in turn, heats the atmosphere
via the processes of conduction, convection (turbulence) and infrared
radiation. As we move upwards from the ground, we might expect temperature
to drop off according to the R-squared law. This happens (more or
less) for a while, but the declining thermal structure reverses at
the tropopause and increases to a new maximum at the stratopause.
In the mesosphere, the temperature drops to the lowest values seen
anywhere in the atmosphere. Above the mesosphere, the temperature
rises again in the thermosphere. Eventually, the temperature reaches
a maximum value at very high altitudes (see Figure above).

Thermal structure of the atmosphere from 0 to 1000 km.
The warmer regions are heated by different parts of the Sun's
radiative output.

The reason for the strange-looking temperature
profile is quite simple. Regions of high temperature are heated
by different portions of the solar radiative output.

The stratosphere is heated by the absorption of ultraviolet (UV)
light by ozone.

The thermosphere is heated by the absorption of extreme ultraviolet
(EUV) light by other atmospheric constituents (primarily molecular
and atomic oxygen and molecular nitrogen).

In this course we are mostly concerned with the troposphere--the
region where we live--and its variations. The stratosphere, however,
also plays a major role in global change and evolution, as we will
soon see. Although some scientists think the mesosphere and thermosphere
might also play key roles in the story of global change, the question
is still not resolved.

We have so far only considered the vertical variation of temperature.
Other atmospheric variables also vary with altitude. Since the atmosphere
is a gaseous envelope, it is compressible. This means that density
and pressure both decrease exponentially with altitude. The lowest
regions are weighed down by the mass of the overlying atmosphere,
becoming compressed and therefore more dense. Jet aircraft flying
in the low-density stratosphere have to pressurize their hulls due
to the compressibility of the atmosphere. Figure 3 shows the variation
of density and pressure with altitude.

The temperature profile shown in Figures 1 and 2 plays a significant
role in controlling atmospheric turbulence. We all know that the troposphere
is a turbulent place to live: we experience wind gusts, cloud formation
and severe weather. The fact that temperature drops with altitude
in the troposphere leads to atmospheric instability. In the stratosphere,
on the other hand, the temperatures rise with altitude, leading to
a very stable region. It is for this very reason that we are able
to drink cups of coffee in jet aircraft. One of the many gifts showered
on us by ozone is the ability to fly commercial aircraft in relative
comfort!

Composition of the Present Atmosphere

The overall composition of the earth's atmosphere is summarized below
along with a comparison to the atmospheres on Venus and Mars - our
closest neighbors.

The variations in concentration from the Earth to Mars and Venus
result from the different processes that influenced the development
of each atmosphere. While Venus is too warm and Mars is too cold for
liquid water the Earth is at just such a distance from the Sun that
water was able to form in all three phases, gaseous, liquid and solid.
Through condensation the water vapor in our atmosphere was removed
over time to form the oceans. Additionally, because carbon dioxide
is slightly soluble in water it too was removed slowly from the atmosphere
leaving the relatively scarce but unreactive nitrogen to build up
to the 78% is holds today.

Current Composition

The concentrations of gases in the earth atmosphere is now known
to be (ignoring water vapor, which varies between near zero to a few
percent):

CONSTITUENT

CHEMICAL SYMBOL

MOLE PERCENT

Nitrogen

N2

78.084

Oxygen

O2

20.947

Argon

Ar

0.934

Carbon Dioxide

CO2

0.035

Neon

Ne

0.00182

Helium

He

0.00052

Methane

CH4

0.00017

Krypton

Kr

0.00011

Hydrogen

H2

0.00005

Nitrous Oxide

N2O

0.00003

Xenon

Xe

0.00001

Ozone

O3

trace
to 0.00080

The unit of percentage listed here are for comparison sake. For most
atmospheric studies the concentration is expressed as parts per million
(by volume). That is, in a million units of air how may units would
be that species. Carbon dioxide has a concentration of about 350 ppm
in the atmosphere (i.e. 0.000350 of the atmosphere or 0.0350 percent).

Global Warming Potential

The Global Warming Potential (GWP) of a greenhouse gas is the ratio
of global warming, or radiative forcing – both direct and indirect –
from one unit mass of a greenhouse gas to that of one unit mass of
carbon dioxide over a period of time. Hence this is a measure of the
potential for global warming per unit mass relative to carbon
dioxide.
Global Warming Potentials are presented for an expanded set of
gases. GWPs are a measure of the relative radiative effect of a
given substance compared to CO2, integrated over a chosen time
horizon. New categories of gases include fluorinated organic
molecules, many of which are ethers that are proposed as halocarbon
substitutes. Some of the GWPs have larger uncertainties than that of
others, particularly for those gases where detailed laboratory data
on lifetimes are not yet available.

The direct GWPs have been calculated relative to CO2
using an improved calculation of the CO2 radiative forcing, the SAR
response function for a CO2 pulse, and new values for the radiative
forcing and lifetimes for a number of halocarbons. Indirect GWPs,
resulting from indirect radiative forcing effects, are also
estimated for some new gases, including carbon monoxide. The direct
GWPs for those species whose lifetimes are well characterized are
estimated to be accurate within ±35%, but the indirect GWPs are less
certain.

Direct Global Warming
Potentials (GWPs) relative to carbon dioxide (for gases for which
the lifetimes have been adequately characterized). GWPs are an index
for estimating relative global warming contribution due to atmospheric
emission of a kg of a particular greenhouse gas compared to emission
of a kg of carbon dioxide. GWPs calculated for different time horizons
show the effects of atmospheric lifetimes of the different gases.

Lifetime

Global
Warming Potential

(years)

(Time
Horizon in Years)

GAS

20
yrs

100
yrs

500
yrs

Carbon Dioxide

CO2

1

1

1

Methane

CH4

12.0

62

23

7

Nitrous Oxide

N2O

114

275

296

156

Chlorofluorocarbons

CFC-11

55

4500

3400

1400

CFC-12

116

7100

7100

4100

CFC-115

550

5500

7000

8500

Hydrofluorocarbons

HFC-23

CHF3

260

9400

12000

10000

HFC-32

CH2F2

5

1800

550

170

HFC-41

CH3F

2.6

330

97

30

HFC-125

CHF2CF3

29

5900

3400

1100

HFC-134

CHF2CHF2

9.6

3200

1100

330

HFC-134a

CH2FCF3

13.8

3300

1300

400

HFC-143

CHF2CH2F

3.4

1100

330

100

HFC-143a

CF3CH3

52

5500

4300

1600

HFC-152

CH2FCH2F

0.5

140

43

13

HFC-152a

CH3CHF2

1.4

410

120

37

HFC-161

CH3CH2F

0.3

40

12

4

HFC-227ea

CF3CHFCF3

33

5600

3500

1100

HFC-236cb

CH2FCF2CF3

13.2

3300

1300

390

HFC-236ea

CHF2CHFCF3

10

3600

1200

390

HFC-236fa

CF3CH2CF3

220

7500

9400

7100

HFC-245ca

CH2FCF2CHF2

5.9

2100

640

200

HFC-245fa

CHF2CH2CF3

7.2

3000

950

300

HFC-365mfc

CF3CH2CF2CH3

9.9

2600

890

280

HFC-43-10mee

CF3CHFCHFCF2CF3

15

3700

1500

470

Fully fluorinated
species

SF6

3200

15100

22200

32400

CF4

50000

3900

5700

8900

C2F6

10000

8000

11900

18000

C3F8

2600

5900

8600

12400

C4F10

2600

5900

8600

12400

c-C4F8

3200

6800

10000

14500

C5F12

4100

6000

8900

13200

C6F14

3200

6100

9000

13200

Ethers and Halogenated
Ethers

CH3OCH3

0.015

1

1

<<1

HFE-125

CF3OCHF2

150

12900

14900

9200

HFE-134

CHF2OCHF2

26.2

10500

6100

2000

HFE-143a

CH3OCF3

4.4

2500

750

230

HCFE-235da2

CF3CHClOCHF2

2.6

1100

340

110

HFE-245fa2

CF3CH2OCHF2

4.4

1900

570

180

HFE-254cb2

CHF2CF2OCH3

0.22

99

30

9

HFE-7100

C4F9OCH3

5

1300

390

120

HFE-7200

C4F9OC2H5

0.77

190

55

17

H-Galden 1040x

CHF2OCF2OC2F4OCHF2

6.3

5900

1800

560

HG-10

CHF2OCF2OCHF2

12.1

7500

2700

850

HG-01

CHF2OCF2CF2OCHF2

6.2

4700

1500

450

Explore the absorption of individual atmospheric gases
and the atmosphere as a whole.

One of the more important factors for climate is the global wind system.
Winds are driven into motion by forces on the air. There is a system of
prevailing winds whose purpose it is to transfer the excess energy received
at low latitudes to high latitudes. If the earth did not rotate and did
not have any continental land masses, then the wind system would be rather
simple.

Figure 6. Wind patterns of the world for (A) a
hypothetical world with no rotation,
(B) the world with rotation, and (C) the resulting bands of general circulation.

The excess heat received in the equatorial region would
cause the air to rise and blow away towards higher latitudes. In order
for air to be conserved, the outward motion at high latitudes near the
equator has to be balanced by inward low altitude winds. This system is
a huge twin-cell circulation pattern (one cell in each atmosphere). These
idealized cells are called Hadley Cells.

Because the earth rotates and has continental land masses, the actual
prevailing winds do not directly blow from pole to equator as in but rather
curve around and form a multicellular circulation pattern. The curving
form the initial direction of the winds is called the "Coriolis effect"
and is due to rotation. The curvature is so great as to split up each
Hadley cell into three smaller cells.

The highest temperatures occur in the subtropical deserts, e.g., the
African Sahara. The lowest mean temperatures occur in Antarctica, where
the Sun is either below the horizon or too low in the sky to effectively
warm the surface.

It is interesting to note here an important feedback process that can
occur at high latitudes. In very cold regions, such as Antarctica, water
can only exist as ice and snow. Both these forms of water are very good
reflectors of visible light (~80% of visible light gets reflected) and
therefore the production of ice and snow tends to depress temperatures
by reflecting much of the potentially warming sunlight back to space!
This is a positive feedback mechanism - cold temperatures lead to ice
- leads to colder temperatures - leads to more ice, etc.

Natural Climate Change

We believe that the temperature of the earth has varied wildly over
the evolution of the earth. Figure 1 shows an estimate of temperature
changes as complied by Scotese. So how can it be that the climate has
changed so over the ages and what processes could lead to these changes?

Figure 1. Estimated changes in global temperature.

The processes for changing climate naturally include:

Plate reorganization

CO2 and Volcanic eruptions

Solar Variations, and

Orbital Variations

Plate Tectonics

The movement of the continents has obviously influenced the climate
at specific locations (Figure 2), but could also influence the global
temperature by redistributing the collection of solar radiation and/or
providing land masses on which continental glaciers could form.

Figure 2. Location of continents during the Devonian
period from Scotese.

Volcanic Eruptions

The amount and location of matgerial added to the atmosphere by volcanoes
probably has significantly influenced climate over the ages. Valcanos
emit some greenhouse gases like CO2 and H2O, but
also emit SO2 that can get trapped in the upper parts of
the atmosphere where it will react to form sulfates, a small particle.
These particles can reflect incoming radiation to lower the surface
temperature.

Figure 3. NOAA monitors the amount or particles (aerosols)
in the atmosphere. Note aerosols over northern and southern Africa.
Click
here
for a current image.

Solar Variability

Variations in sunspot activity result in changes on the order of 0.1%
to 0.2 over 11 year cycle. Numerical climate models predict that a change
in solar output of only 0.5% per century could alter the Earth's climate.

Orbital Variability

The Earth's orbit changes over time in ways that could influence the
amount of energy received at the surface. Thes e include changes in
eccentricity, precession of the equinox, and changes in the Earth's
tilt (obliquity).

Eccentricity

The eccentricity of the Earth's oprbit changes
with a period of 100,000 years. At the moment the Earth's orbit
is fairly circular but in 50,000 years it will be more eccentric
with the difference between aphelion (farthest) and perihelion
(nearest) points in the orbit will become larger.

Precession of the Equinox

Also over time the time of year when the Earth reaches perihelion
changes. Now the Earth is closest the sun in January and faerthest
in July. The combination of changing eccentricity and precession
of the equinox leads to changing available solar radiation.

Obliquity

Finally, the Earth wobbles on its axis of rotation changing
the tilt of the Earth (and hence its seasonality) over a 41,000
year period. The tilt is now 23.5¼ but changes between 22.5¼
and 24.5¼.